Shallow-crustal metamorphism during Late ... - GeoScienceWorld

Shallow-crustal metamorphism during Late Cretaceous anatexis in the Sevier hinterland plateau: Peak temperature conditions from the Grant Range, eastern Nevada, U.S.A. Sean P. Long1*, Emmanuel Soignard2 SCHOOL OF THE ENVIRONMENT, WASHINGTON STATE UNIVERSITY, PULLMAN, WASHINGTON 99164, USA LEROY EYRING CENTER FOR SOLID STATE SCIENCE, ARIZONA STATE UNIVERSITY, TEMPE, ARIZONA 85287, USA

1 2

ABSTRACT Documenting spatio-temporal relationships between the thermal and deformation histories of orogenic systems can elucidate their evolution. In the Sevier hinterland plateau in eastern Nevada, an episode of Late Cretaceous magmatism and metamorphism affected mid- and upper-crustal levels, concurrent with late-stage shortening in the Sevier thrust belt. Here, we present quantitative peak temperature data from the Grant Range, a site of localized, Late Cretaceous granitic magmatism and greenschist facies metamorphism. Twenty-two samples of Cambrian to Pennsylvanian metasedimentary and sedimentary rocks were analyzed, utilizing Raman spectroscopy on carbonaceous material, vitrinite reflectance, and Rock-Eval pyrolysis thermometry. A published reconstruction of Cenozoic extension indicates that the samples span pre-extensional depths of 2.5–9 km. Peak temperatures systematically increase with depth, from ~100 to 300 °C between 2.5 and 4.5 km, ~400 to 500 °C between 5 and 8 km, and ~550 °C at 9 km. The data define a metamorphic field gradient of ~60 °C/km, and are corroborated by quartz recrystallization microstructure and published conodont alteration indices. Metamorphism in the Grant Range is correlated with contemporary, upper-crustal metamorphism and magmatism documented farther east in Nevada, where metamorphic field gradients as high as ~50 °C/km are estimated.These data have implications for localized but significant thermal weakening of the plateau crust, including attaining temperatures for quartz plasticity at depths of ~5–6 km, and the potential for partial melting possibly as shallow as ~12–15 km. Thermal weakening may have contributed to a slowing of shortening rates documented in the Sevier thrust belt at this latitude at ca. 88 Ma, by locally inducing mid- and lower-crustal ductile thickening in the hinterland.

LITHOSPHERE; v. 8; no. 2; p. 150–164; GSA Data Repository Item 2016051 | Published online 25 January 2016

INTRODUCTION

Understanding the thermal history of an orogenic belt, and its relationship in space and time to the corresponding deformation history, can provide critical insights into the evolution of orogenic systems. In the Jurassic-Paleogene U.S. Cordilleran orogenic belt, an episode of Late Cretaceous (ca. 70–90 Ma) granitic magmatism and accompanying metamorphism has been documented in the hinterland plateau of the Sevier thrust belt in eastern Nevada and western Utah (e.g., Miller et al., 1988; Barton, 1990), which produced greenschist and amphibolite facies metasedimentary rocks that are now exposed within a series of highly extended ranges (Fig. 1). However, the complex overprint of Cenozoic extension across Nevada and Utah in many places hinders placing these rocks in an accurate pre-extensional structural and depth context, in particular for greenschist facies rocks that do not contain the mineral assemblages *[email protected]

necessary for quantitative thermobarometry. This, combined with a wealth of studies that have focused on determining the pressure and temperature conditions of upper amphibolite facies rocks exhumed in the footwalls of core complexes (e.g., McGrew et al., 2000; Cooper et al., 2010; Wells et al., 2012; Hallett and Spear, 2014), has led to a prevailing view that metamorphism in the Sevier hinterland dominantly affected mid-crustal levels, despite earlier efforts that presented field relations emphasizing that high temperatures were locally attained in the upper crust (Miller et al., 1988; Miller and Gans, 1989). Part of the difficulty of determining the conditions of shallow-crustal metamorphism is that, until recently, quantitative techniques for analyzing the peak temperature of sub-garnet grade rocks have not been available. Thermal maturation parameters utilized in the petroleum industry, such as vitrinite reflectance, can estimate the peak temperature of sedimentary rocks within ±20–30 °C, but are applicable only up to peak temperatures of ~200–300 °C (e.g., Mukhopad-

doi: 10.1130/L501.1

hyay, 1994). Semiquantitative peak temperature parameters, such as conodont alteration index, only yield broad (~200 °C) bracketing ranges for metasedimentary rocks that experienced peak temperatures ≥300 °C (Königshof, 2003). However, with the development of Raman spectroscopy on carbonaceous material (RSCM) thermometry (Beyssac et al., 2002), peak temperatures for greenschist facies metasedimentary rocks that contain organic matter can be determined within ±50 °C. This technique is ideal for understanding the thermal conditions of thick sedimentary sections such as the Cordilleran passive margin basin in eastern Nevada, which is dominated by organic-rich carbonates (e.g., Stewart, 1980). In this study, we present quantitative peak temperature determinations from Paleozoic sedimentary and metasedimentary rocks in the Grant Range in eastern Nevada, which represents the westernmost of a series of exposures of Late Cretaceous metamorphic rocks in the Sevier hinterland in eastern Nevada (Figs. 1, 2). The samples are placed in a detailed structural

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Shallow-crustal metamorphism in the Sevier hinterland plateau

context, obtained from a recently published structural reconstruction of the Grant Range (Long and Walker, 2015), which shows that they were distributed between crustal depths of ~2.5–9 km prior to Cenozoic extension. Three different thermometry techniques are utilized, including RSCM on Cambrian, Ordovician, and Silurian rocks, and vitrinite reflectance and Rock-Eval pyrolysis on Devonian, Mississippian, and Pennsylvanian rocks. The temperature data are used to quantify a metamorphic field gradient, which is interpreted to reflect the peak thermal conditions for Late Cretaceous metamorphism in the Grant Range. We then compare this data to other studied areas of Late Cretaceous, shallow-crustal metamorphism in eastern Nevada, and discuss implications for localized thermal weakening of the Sevier hinterland plateau crust during the late stages of Cordilleran orogenesis. CORDILLERAN TECTONIC FRAMEWORK OF EASTERN NEVADA

From the Neoproterozoic to the Triassic, >10 km of shallow-marine sedimentary rocks were deposited on the western Laurentian continental shelf in eastern Nevada (e.g., Stewart, 1980). During the Jurassic, establishment of an Andean-style subduction system on the western North American plate margin initiated construction of the Cordilleran orogenic belt, which continued until the Paleogene (e.g., DeCelles, 2004). For the duration of orogenesis, eastern Nevada was situated within a broad hinterland region between the Sierra Nevada magmatic arc and the Sevier fold-thrust belt, the locus of upper-crustal shortening (Fig. 1). The timing of the initiation of shortening in the Sevier thrust belt is debated; most agree that the onset of Early

Sevie rf thrus oldt belt

WUTB

CNTB

c tic ar agma da m Neva Sierra

Figure 1. Map of Cordilleran OR ID 111°W 120°W 114°W 117°W retroarc region in Nevada RMT Salt (NV), Utah (UT), and CaliWY GT Lake fornia (CA) (modified from 41°N Elko City Long, 2015). Deformation fronts of Cordilleran thrust Area of systems shown in blue, Fig. 7 with spatial extents shaded Reno gray. Late Paleozoic thrust LFTB systems shown in brown. Ely 39°N Areas of exposed Jurassic and Cretaceous metamorphic rocks are highlighted, N including metamorphic core complexes in red and Area of Cedar 100 km highly extended ranges Fig. 2A City in orange. AZ—Arizona; UT 37°N AZ CA NV CNTB—Central Nevada thrust belt; GT—Golconda thrust; ID—Idaho; LFTB—Luning-Fencemaker thrust belt; OR—Oregon; RMT—Roberts Mountains thrust; WUTB—Western Utah thrust belt; WY—Wyoming.

Cretaceous (ca. 130 Ma) foreland basin subsidence indicates the onset of shortening (e.g., Jordan, 1981; Currie, 2002), while some argue for emplacement of thrust sheets as early as the latest Jurassic (ca. 145 Ma) (e.g., DeCelles and Coogan, 2006). The onset of crustal thickening in parts of the Sevier hinterland is interpreted to have begun as early as the Late Jurassic (ca. 150 Ma), on the basis of initial prograde metamorphism recorded within rocks now exhumed in the footwalls of metamorphic core complexes (Hoisch et al., 2014; Cruz-Uribe et al., 2015; Kelly et al., 2015) and isotopic compositions of Late Jurassic plutons (Chapman et al., 2015). In Nevada, low-magnitude (a few 10s of km) upper-crustal shortening of likely Cretaceous age was accommodated in the Central Nevada thrust belt (Taylor et al., 2000; Long, 2012; Long et al., 2014) and by regional-scale, open folding across much of eastern Nevada (Long, 2015). However, thermobarometry of exhumed mid-crustal rocks indicates localized but significant Cretaceous crustal thickening, including within the Snake Range core complex in east-central Nevada (e.g., Lewis et al., 1999; Cooper et al., 2010), the Ruby–East Humboldt core complex in northeast Nevada (e.g., Hodges and Walker, 1992; McGrew et al., 2000; Hallett and Spear, 2014), and in the footwall of the Windermere thrust in northeast Nevada (Camilleri and Chamberlain, 1997). By the Late Cretaceous, near the end of crustal thickening, the Sevier hinterland is interpreted to have been a high-elevation orogenic plateau (e.g., Coney and Harms, 1984; DeCelles, 2004), named the “Nevadaplano” after comparison to the Andean Altiplano. The Nevadaplano was underlain by crust that was locally as thick as ~50–60 km (Coney and Harms, 1984; Gans, 1987; DeCelles and Coogan, 2006; Colgan and Henry, 2009;

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Chapman et al., 2015), which supported surface elevations as high as ~3 km (Cassel et al., 2014; Snell et al., 2014). During Cordilleran orogenesis, two episodes of granitic magmatism, one in the Jurassic (ca. 155–165 Ma) and one in the Late Cretaceous (ca. 70–90 Ma), have been documented in eastern Nevada and western Utah, and each is associated spatially and temporally with metamorphism and ductile, contractional deformation (e.g., Barton et al., 1988; Miller et al., 1988; Miller and Hoisch, 1995; Wells and Hoisch, 2008). The Jurassic episode, which in most places is characterized by spatially localized metamorphism in proximity to upper-crustal intrusions (Miller et al., 1988; Miller and Hoisch, 1995), has been interpreted as heating and low-magnitude contractional deformation that accompanied a pulse of back-arc magmatism (Miller and Hoisch, 1995). The more regionally significant Late Cretaceous magmatic event produced the peak metamorphic conditions recorded in much of eastern Nevada (Miller and Gans, 1989; Barton, 1990), and affected both mid- and upper-crustal levels. This event occurred late during the Sevier shortening history, while the Nevadaplano was approaching its maximum crustal thickness (e.g., DeCelles and Coogan, 2006), and was locally associated with significant crustal thickening (Camilleri and Chamberlain, 1997; McGrew et al., 2000). Late Cretaceous magmatism has been interpreted as the shallow thermal expression of lower-crustal anataxis, which was initiated as a result of eastward migration of subduction-related magmatism coupled with conductive relaxation of isotherms within structurally thickened crust (Miller and Gans, 1989), or alternatively, as a result of heat influx following delamination of dense mantle lithosphere (Wells and Hoisch, 2008; Wells et al., 2012). During the Late Cretaceous and Paleocene (ca. 80–60 Ma), spatially isolated extension is recorded in the Nevadaplano, during the final stages of shortening in the Sevier thrust belt (Hodges and Walker, 1992; Camilleri and Chamberlain, 1997; Wells and Hoisch, 2008; Druschke et al., 2009; Long et al., 2015), and has been interpreted as a consequence of thermal weakening and isostatic adjustment following lithospheric delamination (Wells and Hoisch, 2008). In addition, spatially isolated, post-orogenic, Eocene-Oligocene extension has been documented in eastern Nevada (e.g., Gans et al., 2001; Druschke et al., 2009; Long and Walker, 2015), and was often associated in space and time with the sweep of silicic volcanism that accompanied post-Laramide slab rollback (Humphreys, 1995; Dickinson, 2002). The inception of widespread extension that formed

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Eocene-Oligocene dikes Neogene normal faults Paleogene-Neogene detachment faults Mesozoic thrust faults and folds Drill holes that intercept granite

Eocene-Oligocene volcanic and sedimentary rocks Jurassic-Cretaceous granite Mississippian-Pennsylvanian sedimentary rocks Silurian-Devonian sedimentary rocks Cambrian-Ordovician sedimentary rocks (hashes indicate metamorphism)

Figure 2. (A) Map showing location of the Grant Range and geographic names of surrounding ranges and valleys. (B) Generalized geologic map of the Grant Range, modified from Long and Walker (2015). Compiled source maps include Moores et al. (1968), Kleinhampl and Ziony (1985), Lund et al. (1987; 1988), Fryxell (1988), Camilleri (2013), and Long (2014). Dashed pink line shows outer limit of positive aeromagnetic anomaly interpreted to delineate subsurface extent of Troy pluton (Lund et al., 1987, 1988; Blank, 1993). Locations of Railroad Valley drill holes that intercept granite at depth are from Hess et al. (2004).

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the Basin and Range province, and accompanying lowering of the surface elevation of the Nevadaplano, was not until the middle Miocene (e.g., Dickinson, 2002; Colgan and Henry, 2009; Cassel et al., 2014). DEFORMATION, METAMORPHISM, AND MAGMATISM IN THE GRANT RANGE

An 8-km-thick section of Cambrian-Pennsylvanian sedimentary rocks is exposed in the Grant Range (Figs. 2, 3), and is dominated by carbonates (Fryxell, 1988; Lund et al., 1993; Camilleri, 2013; Long and Walker, 2015). Cambrian and Ordovician rocks have experienced greenschist facies metamorphism and low-magnitude, penetrative contractional strain (Figs. 4B, 4D), both of which die out upsection, and are not observed in Silurian-Pennsylvanian rocks (Fryxell, 1988; Lund et al., 1993; Camilleri, 2013). Readers are referred to Camilleri (2013) for detailed descriptions of micro- and mesoscale ductile structures and metamorphic lithologies, textures, and mineral assemblages in the study area in the central Grant Range (Fig. 2). Rock unit divisions discussed below and shown on Figure 3 are after Long and Walker (2015). Cambrian and Ordovician sandy and silty limestone protoliths have been metamorphosed to recrystallized limestone with abundant mica porphyroblasts, with silty partings metamorphosed to argillite and phyllite (Figs. 4A, 4C), and more pure limestone protoliths have been metamorphosed to crystalline marble (Figs. 4B, 4D). Moving downsection through the stratigraphic column (Fig. 3), sericite (defined here as fine-grained white mica) and chlorite first appear within upper Ordovician rock units, tourmaline and white mica porphyroblasts appear at the top of the Cambrian section, and phlogopite, amphibole (Fig. 4C), and biotite porphyroblasts are observed within the lowest Cambrian units in the study area (Camilleri, 2013). All metasedimentary rock units observed in the study area have limestone protoliths, and therefore variation in protolith composition is not interpreted to have strongly affected the resulting mineral assemblages. In the southern Grant Range, stratigraphically lower Cambrian rocks are exposed, including the Pioche shale and Prospect Mountain quartzite (Fig. 3). Both of these units exhibit biotite porphyroblasts, and in one locality ~5 km south of the map area, the Pioche shale exhibits staurolite porpyhroblasts (Fryxell, 1988). Three sequential phases of metamorphism are recorded in Cambrian and Ordovician rocks in the study area (Camilleri, 2013). The first phase accompanied mesoscale, east-vergent folding, and is defined by growth of oriented

white mica porphyroblasts during development of axial-planar cleavage. The second phase involved growth of randomly oriented white mica, phlogopite, biotite, and amphibole porphyroblasts, which are interpreted as static metamorphic textures. On the basis of mineral assemblage, this stage is interpreted to have yielded the peak metamorphic conditions recorded in the Grant Range (Camilleri, 2013). This was followed by an additional stage of synkinematic metamorphism that accompanied west-vergent, mesoscale folding and thrust faulting, and is characterized by growth of oriented white mica and chlorite during development of axial-planar cleavage (Camilleri, 2013). In the southern Grant Range, 3–10 km south of the study area, Cambrian rocks record a history of metamorphism and large-scale folding, and are intruded by the Troy granite stock (Fryxell, 1988) (Fig. 2). U-Pb zircon geochronology indicates two distinct emplacement ages for the stock (Lund et al., 2014). A boudinaged granite sill on the western margin of the stock, which pre-dates folding, yielded a Jurassic (163.3 ± 1.2 Ma) age, and undeformed granite that makes up the bulk of the pluton, and post-dates folding, yielded a Late Cretaceous age (83.7 ± 0.8 Ma) (Lund et al., 2014). A stage of static, peak metamorphism, indicated by randomly oriented porphyroblasts, is recorded in the southern Grant Range, and post-dates large-scale folding (Fryxell, 1988). The metamorphic grade of Cambrian sedimentary rocks generally decreases upsection in the southern Grant Range, but Fryxell (1988) also documented that metamorphism in Cambrian rocks dies out to the south, with distance from the granite. This suggests that Late Cretaceous magmatism may have been a primary source of heat for metamorphism. After these observations in the southern part of the range, the stage of static, peak metamorphism in the study area in the central Grant Range has also been interpreted to have been, at least in part, contemporary with intrusion of the Late Cretaceous component of the Troy stock (Lund et al., 1993; Camilleri, 2013). This interpretation is supported by geophysical and drill hole data that suggest that the Troy pluton extends in the subsurface under the west-central Grant Range. A positive aeromagnetic anomaly, interpreted to delineate the subsurface extent of the Troy pluton (Lund et al., 1987, 1988; Blank, 1993), extends through the western part of the study area (Fig. 2). Several drill holes in Railroad Valley that fall within the area of the aeromagnetic anomaly intercept granite at depth (Fig. 2) (Lund et al., 1993; Hess et al., 2004; Long and Walker, 2015), which supports this interpretation. In the southernmost Grant Range and farther south in the Quinn Canyon Range, east-vergent


thrust faults correlated with the Central Nevada thrust belt deform Cambrian through Devonian rocks (Taylor et al., 2000) (Fig. 2). However, in the central and northern Grant Range, regionally continuous, older-over-younger relationships indicative of large-scale thrust faults have not been documented through the exposed thickness of Cambrian to Pennsylvanian rocks (Lund et al. 1993; Camilleri, 2013; Long and Walker, 2015). Paleozoic rocks in the Grant Range are unconformably overlain by Paleogene sedimentary and volcanic rocks. 1300 m of Paleogene rocks are preserved in the study area (Long and Walker, 2015) (Fig. 3), but sections as thick as ~2000 m are reported in other parts of the range (Moores et al., 1968). In the study area, Paleogene rocks overlie Mississippian and Pennsylvanian rock units, with minimal angular discordance (Long and Walker, 2015). In addition, Eocene-Oligocene andesite, dacite, and granite dikes intruded Cambrian and Ordovician stratigraphic levels (Fig. 2) (Fryxell, 1988; Camilleri, 2013; Long and Walker, 2015). In the study area, a dacite dike yielded a crystallization age of ca. 29 Ma (40Ar/39Ar biotite; Long and Walker, 2015), and in the southern Grant Range, a granite dike is dated at 31.7 ± 0.8 Ma (U-Pb zircon; Lund et al., 2014). Paleozoic and Paleogene rocks in the Grant Range experienced a polyphase history of Cenozoic extension. The earliest extension was accommodated by low dip-angle, top-to-west detachment faults (Lund et al., 1993; Camilleri, 2013; Long and Walker, 2015), which initiated in the Oligocene (ca. 29–32 Ma) (Long and Walker, 2015). This was followed by one or more episodes of extension accommodated by high-angle normal faults, including Miocene to Quaternary faulting associated with formation of the Railroad Valley structural basin (Fig. 2) (Moores et al., 1968; Lund et al., 1993; Camilleri, 2013). PEAK TEMPERATURE DATA Structural, Stratigraphic, and Depth Context of Samples

A suite of 22 samples of Paleozoic sedimentary rocks ranging in age from Cambrian to Pennsylvanian were analyzed for the peak temperature that they experienced, using three different thermometry techniques, RSCM, vitrinite reflectance, and Rock-Eval pyrolysis. All samples were collected in the central Grant Range, within the map area of Long and Walker (2015) (Figs. 2, 5A). Samples are given structural context by projecting them to their present-day positions on the deformed cross-

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0˚

Pennsylvanian

Pennsylvanian rocks, undifferentiated (Egan Range)

1,000

Goodwin Formation

Stratigraphic depth (m) 5,000 4,000

Little Meadows Fm. Blue Eagle member

6,000 7,000

Willow Spring member Pole Canyon Limestone

10,000

Prospect Mountain Quartzite

Cpm

59.5°C / km

Mineral appearance isograds from Camilleri (2013)

sericite-in

Peak temperature data

projected paleo-surface level (15 ± 10 °C) Rock-Eval pyrolysis (±0.2 VRo, n=3) vitrinite reflectance (2σ, n=6)

GR50 white mica-in, GR67 tourmaline-in

Raman spectroscopy on carbonaceous material (2 SE, n=16)

GR69

subgrain rotation (SGR) quartz recrystallization (ca. 400-500 °C, n=2)

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Best-fit metamorphic field gradient

3-5 CAI (n=12)

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marble

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Opl Opk Ops GR35 Opp GR34 chlorite-in

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Guilmette Formation

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y = 16.798x - 250 R 2= 0.8547

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Lithologies

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Simonson Dolomite Oxyoke Canyon Sandstone Sevy Dolomite Laketown Dolomite Ely Springs Dolomite Eureka Quartzite Lehman Formation Kanosh Shale Shingle Limestone

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Peak Temperature (˚C) 100˚ 200˚ 300˚ 400˚ 500˚ 600˚ 700˚

GR58 phlogopite-in, amphibole-in GR55 GR53 GR56 GR64 biotite-in GR59 GR49

general range from conodont alteration index (CAI) of rocks within 75 km radius (Crafford, 2007) metamorphic temperature range for biotite-bearing units in southern Grant Range (Fryxell, 1988) metamorphic temperature range for staurolite-bearing Cp in southern Grant Range (Fryxell, 1988)

base not exposed

Figure 3. Stratigraphic column of Grant Range rock units, using divisions and thicknesses from Long and Walker (2015) (lithologies and estimated eroded thicknesses of Pennsylvanian and Permian rocks in the Egan Range are from Kellogg, 1963, and Brokaw and Heidrick, 1966). Column on right shows peak temperature versus stratigraphic depth below projected paleo-surface level for the 22 analyzed samples.


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B

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Figure 4. Photographs and photomicrographs, illustrating: (A) Thin-bedded, organic-rich limestone with phyllitic partings, characteristic of unit _sgl (refer to Fig. 3 for a guide to all unit abbreviations used in this caption). (B) Marble characteristic of unit _sw, with bedding (subhorizontal) overprinted by spaced cleavage (dipping toward left-hand side). (C) Amphibole (black) and white mica (silver) porphyroblasts within a phyllitic parting in marble of unit _sw. (D) West-vergent, mesoscale fold in phyllitic marble of unit _pc. (E) Black microlithons of carbonaceous material (CM) in marble of unit _pc (sample GR64; plane-polarized light). (F) Equigranular, polygonal quartz subgrain microstructure characteristic of subgrain rotation (SGR) recrystallization (sample GR67; cross-polarized light). (G, H, I) Representative examples of analyzed CM (G: sample GR50, spot 9; H: sample GR36, spot 3; I: sample GR55, spot 6; all taken in plane-polarized light; green spot in each photo is the Raman laser beam).


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bedrock culmination axis GR38 peak temperature sample

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Figure 5. (A) Geologic map of the central Grant Range, generalized from Long and Walker (2015). (B, C) Cross-sections A–A′ and B–B′, generalized from Long and Walker (2015), showing present-day geometry. Translucent areas above modern erosion surface represent eroded rock. Peak temperature samples are projected along-strike to their sampled structural level. (D, E) Cross sections A–A′ and B–B′, restored for motion on Cenozoic normal and detachment faults and for the magnitude of flexural isostatic folding that accompanied extension; generalized from Long and Walker (2015). Translucent areas represent rock that has either been eroded above the modern erosion surface or translated westward off of the map area during extension. The restored positions of peak temperature samples are projected.


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sections of Long and Walker (2015) (Figs. 5B, 5C). In addition, Long and Walker (2015) restored Cenozoic extension in the Grant Range by retro-deforming offset on all normal and detachment faults, and presented observations indicating that Paleozoic rocks had gentle dip magnitudes prior to extension, including documenting minimal angular discordance and structural relief across the Paleogene unconformity, and retro-deformation of the magnitude of flexural isostatic folding that accompanied extension. The samples are projected to their pre-extensional positions on the restored crosssections of Long and Walker (2015) (Figs. 5D, 5E), which reveals that they were spread over a total east-west distance of 30–40 km prior to extension. Based on the reconstruction of Long and Walker (2015), we assume that stratigraphic depths for the samples are approximately equivalent to pre-extensional structural depths. Samples are shown in a stratigraphic column (Fig. 3) using unit thicknesses from Long and Walker (2015), which are based on a combination of complete unit thicknesses measured in their cross sections, complete thicknesses reported in other published studies in the Grant Range (Moores et al., 1968; Fryxell, 1988; Camilleri, 2013), and minimum tectonic thicknesses for some Cambrian and Ordovician rock units. Despite using minimum thicknesses, the cumulative thicknesses of the Cambrian and Ordovician sections in the study area are comparable to nearby estimates in the White Pine and Egan Ranges (Kellogg, 1963; Moores et al., 1968), and regionally in eastern Nevada (Stewart, 1980).

The samples span a total stratigraphic thickness of 6.5 km (Fig. 3). However, assigning preextensional stratigraphic depths is more difficult, as an unknown thickness of rocks was eroded off of the study area prior to the Paleogene (e.g., Long, 2012). In the central Grant Range, only a few hundred meters of Pennsylvanian rocks are preserved beneath the Paleogene unconformity (Fig. 3). However, incomplete and complexly faulted sections of Pennsylvanian and Permian rocks are preserved along-strike to the north in the White Pine Range (Moores et al., 1968), which suggests that stratigraphic levels at least as high as Permian were at one time present over the study area. To the east in the Egan Range, Kellogg (1963) and Brokaw and Heidrick (1966) documented a total thickness of 2.0–2.5 km of Pennsylvanian and Permian rocks. Conodont alteration indices from Pennsylvanian and Permian rocks within a ~75 km radius of the study area are characterized by values of 1–1.5 (n = 19) (Crafford, 2007), which corresponds to a maximum burial temperature range of 50–80 °C (Königshof, 2003). These data indicate that Pennsylvanian and Permian rocks in the study area were not deeply buried by a thick Mesozoic section that is now eroded away (e.g., Long, 2012). On the basis of westward onlap of Triassic rocks and westward erosional truncation of Permian rocks below Triassic rocks, several studies have argued that east-central Nevada (including the study area) was a topographic high during much of the Triassic (Burchfiel et al., 1974; Collinson et al., 1976; Stewart, 1980), and therefore did not accumulate a thick section of Triassic rocks. Therefore, samples are reported

on Figure 3 and Tables 1–3 at their stratigraphic depth below 2.5 km of eroded Pennsylvanian and Permian rocks, after thicknesses reported in the Egan Range (Kellogg, 1963; Brokaw and Heidrick, 1966). Based on the regional observations summarized above, these stratigraphic depths are interpreted to represent maximum permissible estimates for burial depths of rocks in the study area prior to synorogenic erosion. RSCM Thermometry

Carbonaceous material (CM), which is derived from solid-state metamorphism of organic material (e.g., Buseck and Huang, 1985), is common in many metasedimentary rocks. Several studies have shown that the degree of structural organization of graphite bonds in CM is strongly temperature dependent, and therefore can be used as a quantitative geothermometer (e.g., Beyssac et al., 2002; Rahl et al., 2005; Aoya et al., 2010). Rahl et al. (2005) calibrated the RSCM thermometer with an uncertainty of ± 50 °C (2s) for rocks that achieved peak temperatures between 100 and 700 °C, by measuring the height ratio (R1) and area ratio (R2) of four first-order Raman peaks (G, D1, D2, D3) in the wavenumber offset range between 1200 and 1800 cm–1. Here, R1, R2, and peak temperature are determined from Equations 1, 2, and 3, respectively, of Rahl et al. (2005). CM was analyzed from 16 samples of Cambrian, Ordovician, and Silurian rocks, with lithologies including marble, phyllite, limestone, and dolomite (Table 1; Fig. 3). Most samples contained abundant CM, which was either pres-

TABLE 1. PEAK TEMPERATURE DETERMINATIONS FROM RAMAN SPECTROSCOPY ON CARBONACEOUS MATERIAL Sample

GR29 GR38 GR36 GR35 GR34 GR50 GR67 GR69 GR68 GR58 GR55 GR53 GR56 GR64 GR59 GR49

Map unit

Lithology

Sl Sl Sl Ops Opp _sb _sgu _sgu _sgl _sw _sw _sw _pc _pc _pc _pc

dolomite dolomite dolomite limestone limestone limestone limestone limestone limestone marble phyllite marble marble marble marble marble

Stratigraphic depth (m)

Latitude (dd.ddddd)

Longitude (dd.ddddd)

4075 4090 4210 5095 5235 6150 6325 6855 8090 8470 8820 8910 9005 9005 9050 9080

38.44536 38.44267 38.44267 38.44250 38.44367 38.44667 38.44711 38.43297 38.42447 38.45214 38.45019 38.43944 38.45050 38.45531 38.46950 38.44903

115.42425 115.42975 115.42975 115.43325 115.43739 115.45867 115.48631 115.51264 115.52911 115.48531 115.47253 115.47581 115.45886 115.46428 115.47017 115.46467

R1

R2

Peak temperature (°C)

Mean

1σ

Mean

1σ

Mean

1σ

2 SE

1.466 1.657 1.984 0.245 0.261 0.190 0.309 0.463 0.303 0.285 0.139 0.117 0.161 0.205 0.176 0.211

0.224 0.112 0.485 0.028 0.161 0.098 0.083 0.131 0.065 0.123 0.055 0.043 0.101 0.045 0.046 0.064

0.721 0.7704 0.766 0.349 0.336 0.335 0.348 0.414 0.330 0.338 0.221 0.176 0.224 0.283 0.270 0.257

0.043 0.026 0.071 0.045 0.100 0.111 0.057 0.070 0.051 0.087 0.079 0.054 0.100 0.061 0.043 0.067

261 225 221 438 455 437 457 425 475 461 544 586 547 498 503 527

42 26 80 44 67 95 43 46 42 62 69 46 84 53 35 57

36 30 50 37 48 68 34 36 36 44 54 41 51 40 32 41

n

13 14 14 13 12 10 15 14 13 13 10 11 15 13 14 14

Note: (1) R1, R2, and peak temperature values calculated using the calibration of Rahl et al. (2005). Internal variability in R1, R2, and peak temperature is indicated by 1σ uncertainty. Temperature is also reported with 2 standard errors (SE), calculated after Cooper et al. (2013), from quadratic addition of 1σ internal error and external error of ±50 °C from the Rahl et al. (2005) calibration, divided by the square root of the number of analyses (n). (2) Stratigraphic depths calculated below estimated paleo-surface level at 2.5 km above base of unit IPe, from thicknesses of Pennsylvanian and Permian rocks in the Egan Range (Kellogg, 1963; Brokaw and Heidrick, 1966).


157

LONG AND SOIGNARD

ent within organic-rich microlithons (Figs. 4E, 4I), or as isolated patches, typically ≤50 µm in diameter (Fig. 4G, 4H). CM was analyzed in situ on polished petrographic thin sections that were cut normal to bedding. Measurements were made at the LeRoy Eyring Center for Solid State Science at Arizona State University, using a custom-built Raman spectrometer. The 532 nm laser was operated at a power of 3 mW, and was focused using a 50× ultra-long working distance Mitutoyo objective. The probed area of CM for each measurement was ~1 µm in diameter (Figs. 4G, 4I). Instrument parameters, settings, and procedures follow those outlined in Cooper et al. (2013). Where possible, the laser was focused on CM situated beneath a transparent grain (typically calcite), after procedures outlined in Beyssac et al. (2003). CM was analyzed for 120 seconds over a spectral window of 1100–2000 cm–1, and typically 15 separate spots were analyzed in each sample, to allow evaluation of in-sample variation. Peak positions, heights, widths, and areas of the Raman spectra (see supplementary information for supporting data from individual analyses1) were determined using a custom peak fitting program written in Matlab by E. Soignard. The program allowed peak shapes to be fit by a combination of gaussian and lorentzian peaks, and any background slope to be removed by using a first-order polynomial between 1100 and 2000 cm–1. Examples of representative Raman spectra for each sample are shown in Figure 6, and summary data for peak temperature determinations are shown in Table 1. Mean peak temperatures of multiple measurements are reported on Table 1, and internal uncertainty within each sample is represented by the reported 1s error in R1, R2, and peak temperature. However, after Cooper et al. (2013), peak temperatures are reported with 2 standard errors (SE), which takes into account the external error of ± 50 °C from the Rahl et al. (2005) calibration (see footnote of Table 1). At 2 SE, typical error ranges are ± 30–50 °C (Table 1). Resulting RSCM peak temperatures exhibit a total range between ~260 and 580 °C (Table 1; Fig. 3), and temperatures generally increase with stratigraphic depth. The three Silurian samples (GR29, GR38, GR36) yielded temperatures between ~220 and 260 °C. Seven samples from the Ordovician section and the upper and middle parts of the Cambrian section (GR35, GR34, GR50, GR67, GR69, GR68, GR58) range between ~420 and 480 °C. The lowest six samples (GR55, GR53, GR56, GR64, GR59, GR49), 1 GSA Data Repository Item 2016051, Tables DR1– DR3, Figure DR1 and related supporting information, is available at www.geosociety.org/pubs/ft2016.htm, or on request from [email protected].

from Cambrian units _pc and _sw (see Figure 3 for a guide to all unit abbreviations used in the text), range between ~500 and 580 °C.

GR29, spot 10 T=246°C

R1=1.653 R2=0.751

Vitrinite Reflectance Thermometry

GR38, spot 14 T=224°C

R1=1.407 R2=0.755

GR36, spot 7 T=234°C

R1=1.678 R2=0.764

Vitrinite reflectance, a thermal maturation parameter widely applied in the petroleum industry, can be used to estimate the peak temperature of sedimentary rocks over the range of ~50–300 °C (e.g., Barker and Pawlewicz, 1994; Mukhopadhyay, 1994). Random vitrinite reflectance, the proportion of normal incident light reflected by a polished surface of vitrinite, increases systematically with peak temperature, as the result of a series of chemical transformations that accompany hydrocarbon maturation (Mukhopadhyay, 1994). Six samples of limestone, dolomite, and shale from Devonian, Mississippian, and Pennsylvanian rocks were analyzed for vitrinite reflectance thermometry (Table 2). Analyses were performed by Weatherford Laboratories, Inc., using procedures outlined in ASTM (2014). No primary vitrinite fragments were identified in the samples; instead, measurements of random reflectance were made on grains of solid bitumen, a vitrinite-like maceral (e.g., Landis and Castano, 1995). As few as two and as many as 30 measurements of solid bitumen reflectance (RSB) were made from individual samples (supporting data in supplementary information; see footnote 1). RSB values were converted into equivalent vitrinite reflectance (RVE) values using the equation of Jacob (1989) (Table 2). Peak temperatures were then obtained from the mean RVE value, using the calibration equation for burial heating from Barker and Pawlewicz (1994), as samples were not collected in stratigraphic continuity to allow proper thermal modeling. Error is reported at the 2s level, which corresponds to a typical error range of ±10–25 °C. Peak temperatures determined from the six samples range between ~130 and 150 °C (Table 2; Fig. 3), and temperatures from all samples overlap within estimated error. Rock-Eval Pyrolysis Thermometry

Rock-Eval pyrolysis, another technique widely applied in the petroleum industry to estimate the thermal maturity of organic matter, can be used to estimate the peak temperature of sedimentary rocks over the range of ~50–200 °C (e.g., Clementz, 1979; Peters, 1986). Pyrolysis involves heating of a pulverized rock sample in the absence of oxygen, in order to release and measure organic compounds (Peters, 1986). Tmax, the oven temperature at which the maximum amount of non-volatile hydrocarbons are released, can be converted into a calculated vi-


R1=0.229 R2=0.341

GR34, spot 19 T=456°C

R1=0.180 R2=0.315


R1=0.385 R2=0.463


R1=0.388 R2=0.372

GR69, spot 17 T=424°C

R1=0.491 R2=0.423

GR68, spot 10 T=473°C

R1=0.285 R2=0.327

GR58, spot 12 T=486°C

R1=0.212 R2=0.296

GR55, spot 13 T=544°C

R1=0.119 R2=0.216


R1=0.148 R2=0.182


R1=0.177 R2=0.239


R1=0.222 R2=0.284

GR59, spot 14 T=506°C

R1=0.136 R2=0.256

GR49, spot 10 T=539°C D1

G

R1=0.198 R2=0.242 D2

1200 1300 1400 1500 1600 1700 1800 Raman Shift (cm-1)

Figure 6. Examples of representative Raman spectra from each of the 16 samples analyzed for Raman spectroscopy on carbonaceous material. The positions of the graphite band (G) and defect bands (D1, D2) are shown on the bottom spectrum. Samples are stacked in stratigraphic order. Peak temperatures (T) and R1 and R2 parameters are calculated after Rahl et al. (2005). Peak center position, height, amplitude, and area are listed for individual analyses in the supplementary information (see footnote 1).


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TABLE 2. PEAK TEMPERATURE DETERMINATIONS FROM VITRINITE REFLECTANCE Sample

Map unit

Lithology

Stratigraphic depth (m)

Latitude (dd.ddddd)

Longitude (dd.ddddd)

Mean RSB (%)

RSB (1σ)

Mean RVE (%)

RVE (1σ)

n

Tpeak (°C)

± 2σ error (°C)

GR09 GR08B GR14 GR04 GR42 GR37

IPe Mc Mc Mc Dsi Dse

limestone siltstone shale limestone dolomite dolomite

2435 2695 2845 2860 3875 3975

38.46167 38.46536 38.45800 38.47006 38.45653 38.42458

115.37467 115.38319 115.40414 115.51256 115.43808 115.42819

1.14 1.17 0.95 1.26 0.91 0.91

0.253 0.368 0.150 0.136 0.117 0.196

1.11 1.12 0.99 1.18 0.96 0.96

0.146 0.161 0.090 0.083 0.071 0.112

8 2 16 30 23 7

144 145 135 149 132 132

+19/–25 +20/–27 +13/–16 +11/–12 +11/–13 +21/–17

Note: (1) RSB was converted into RVE using the equation of Jacob (1989): RVE = 0.618*RSB + 0.40. (2) Tpeak calculated from RVE values, using equation for burial heating from Barker and Pawlewicz (1994): Tpeak = (ln(RVE) + 1.68)/0.0124. (3) Stratigraphic depths calculated below estimated paleosurface level at 2.5 km above base of unit IPe, from thicknesses of Pennsylvanian and Permian rocks in the Egan Range (Kellogg, 1963; Brokaw and Heidrick, 1966). Abbreviations: RSB = solid bitumen reflectance, RVE = equivalent vitrinite reflectance (from Jacob, 1989), Tpeak = peak temperature.

trinite reflectance value (cal. RV) (Jarvie et al., 2001), and therefore can be used to estimate peak temperature. Three samples of limestone, siltstone, and shale from Mississippian and Pennsylvanian rock units were analyzed for Rock-Eval pyrolysis (Table 3). Analyses were performed by Weatherford Laboratories, Inc., using instrument settings and procedures outlined in Clementz (1979) and Peters (1986). Additional supporting data are included in the supplementary information (see footnote 1). Tmax was converted into cal. RV using the equation of Jarvie et al. (2001), and peak temperature was calculated from cal. RV using the calibration of Barker and Pawlewicz (1994). Peak temperatures for the three samples range between ~100 and 125 °C (Table 3). In the absence of a statistically meaningful procedure for calculating uncertainty, errors of ±0.2 cal. RV were assigned, which results in a typical temperature error range of ±15–30 °C, which is comparable to the ± 2s error range for the vitrinite reflectance analyses. Peak temperatures for all three pyrolysis samples overlap within this estimated error range (Table 3). Peak temperature estimates for these three samples were obtained from both vitrinite reflectance and pyrolysis. Pyrolysis temperature estimates were consistently lower, by a margin of ~20–45 °C; however, temperatures from both techniques overlap within estimated error for all three samples.

Additional Peak Temperature Constraints

Three additional data sets, including quartz recrystallization microstructure, published conodont alteration indices, and published metamorphic temperature ranges based on mineral assemblage, supply semiquantitative peak temperature constraints for Paleozoic rocks in the Grant Range. Characterization of the morphology of dynamic recrystallization of quartz in thin section allows bracketing of deformation temperature range (e.g., Stipp et al., 2002). In the study area, two samples of sandy limestone from map units _sgu (GR67) and _sb (GR50) display evidence for subgrain rotation recrystallization (Fig. 4F), which is characterized by ~0.1 mm, equigranular, polygonal quartz subgrains, and indicates deformation temperatures of ~400–500 °C (Stipp et al., 2002) (Fig. 3). This temperature range is in agreement with RSCM peak temperatures from these samples (457 ± 34 °C, 437 ± 68 °C). Conodont alteration indices (CAI) provide semiquantitative estimates of the peak temperature that sedimentary rocks have experienced during diagenetic burial or metamorphism (e.g., Epstein et al., 1977; Königshof, 2003). Compilation of characteristic CAI values for Ordovician through Pennsylvanian sedimentary rocks within a 75 km radius of the study area, using the database of Crafford (2007), gives a general range of peak temperature versus stratigraphic level for

the upper 6 km of the Paleozoic section (Fig. 3). Pennsylvanian-Permian rocks are characterized by CAI values of 1–1.5 (